porosity evolution and mineral paragenesis during low-grade

POROSITY EVOLUTION AND MINERAL PARAGENESIS DURINGLOW-GRADE METAMORPHISM OF BASALTIC LAVAS AT

TEIGARHORN, EASTERN ICELAND

PHILIP S. NEUHOFF*, THRAINN FRIDRIKSSON*, STEFAN ARNORSSON**,and DENNIS K. BIRD*

ABSTRACT. Low-grade alteration of basaltic lavas at Teigarhorn, easternIceland, resulted in three distinct stages of mineral paragenesis that correlate toevents in the burial and intrusive history of Icelandic crust. Metasomatism andbrittle deformation during the paragenetic stages dramatically affected the paleo-hydrology of the lavas and formed temporally distinct mineral assemblages. Thelavas initially contained up to 22 percent total porosity concentrated near thetops and bottoms of individual lava flows. Celadonite and silica (Stage I) precipi-tated along the walls of primary pores prior to deep burial of the lavas andoccluded �8 percent of the initial porosity. During burial (Stage II), hydrolysis ofolivine and basaltic glass led to the formation of mixed-layer chlorite/smectiteclays in the matrix of the lavas and as rims filling �40 percent of the volume ofprimary pores. Chlorite contents in Stage II mixed-layer clay rims increased from�20 to �80 percent during the infilling of individual vesicles, reflecting increas-ing temperatures with time as the lavas were buried. The end of Stage II occurredafter burial and is represented by filling of �40 percent of total primary porosityby zeolites (scolecite or heulandite � stilbite � mordenite � epistilbite) and re-placement of plagioclase by zeolites and albite. The Stage II zeolite assemblagesare indicative of two regional metamorphic mineral zones in eastern Iceland, themesolite � scolecite and heulandite � stilbite zones. The presence of the bound-ary between the mesolite � scolecite � and heulandite � stilbite zones indicatesthat the maximum temperature during burial metamorphism was 90° � 10°C.Localized areas of intense hydrothermal alteration associated with intrusion ofbasaltic dikes (Stage III) overprint Stages I and II. Extensive fracturing andhydrothermal brecciation during Stage III added 3 to 11 percent total porosity inwhich mm- to cm-scale museum-grade crystals of quartz, calcite (Iceland spar),stilbite, scolecite, heulandite, and laumontite precipitated. Estimates of the tem-perature during Stage III (based on fluid inclusion homogenization temperaturesin calcite, chlorite geothermometry, and the zeolite assemblage) range from 120°to 180°C. Although the thermobarometric conditions during Stages II and III ledto similar mineral assemblages, careful attention to textural and geologic rela-tions permits seemingly complex, multi-stage parageneses in metabasalts to beinterpreted in a petrotectonic context.

INTRODUCTION

The mineralogy of low-grade metabasalts is a sensitive indicator of the thermobaro-metric evolution of oceanic crust and continental flood basalts. Discrete temperature-,pressure-, and/or depth-controlled zones characterized by assemblages of silica miner-als, phyllosilicates, and zeolites frequently serve as metamorphic indicators in low-grademetabasalts. Zones containing trioctahedral smectite, corrensite and/or mixed layerchlorite/smectite, and chlorite occur with increasing metamorphic grade in zeolite- andgreenschist-facies metabasalts and active geothermal systems in basaltic rocks (Mehegan,Robinson, and Delany, 1982; Seki and others, 1983; Liou and others, 1985; Bevins,Robinson, and Rowbotham, 1991; Robinson, Bevins, and Rowbotham, 1993; Schmidt,1990; Schmidt and Robinson, 1997; Kristmannsdottir, 1979; Schiffman and Fridleifsson,1991; Robinson and Bevins, 1999; Alt, 1999). In addition, as many as five separatedepth-controlled ‘‘zeolite zones’’ have been described in regionally metamorphosedbasaltic lavas and Icelandic geothermal systems (Walker, 1951, 1960a,b; Sukheswala,

* Department of Geological and Environmental Science, Stanford University, Stanford, California94305-2115

** Science Institute, University of Iceland, Dunhagi 3, 107 Reykjavik, Iceland

[AMERICAN JOURNAL OF SCIENCE, VOL. 299, JUNE, 1999, P. 467–501]

467

Avasia, and Gangopadhyay, 1974; Kristmannsdottir and Tomasson, 1978; Jørgensen,1984; Murata, Formoso, and Roisenberg, 1987; Schmidt, 1990; Neuhoff and others,1997; Christiansen and others, 1999). Phyllosilicate and zeolite parageneses thus provideuseful information for evaluating heat flow and structural evolution in large igneousprovinces (Neuhoff and others, 1997) and exploration potential in fractured basalticpetroleum reservoirs (Christiansen and others, 1999). Complicating such interpretations,however, is the fact that basaltic provinces often undergo multiple stages of intrusion,extrusion, and deformation during which the lavas are altered through weathering,regional burial metamorphism, and local hydrothermal/contact metamorphism. Min-eral assemblages formed during regional metamorphism and localized hydrothermalalteration can be nearly identical, complicating interpretation of mineral isogradsassociated with burial metamorphism (Neuhoff and others, 1997; Sukheswala, Avasia,and Gangopadhyay, 1974).

In the present study, we demonstrate that regional and local mineral parageneses inlow-grade metabasalts can be distinguished on the basis of geologic and texturalrelations. These relations are difficult to establish from drillcores and areas of limitedoutcrop. Ideally one would like to investigate the textural and geologic relations betweenmineral parageneses in laterally continuous, unweathered outcrop exposures. Remark-able outcrop exposures created through coastal erosion at Teigarhorn, eastern Iceland(fig. 1), permit such detailed observations of geologic and textural relationships inbasaltic lavas affected by burial metamorphism, brittle deformation, and hydrothermalalteration during dike emplacement. Geologic, petrographic, mineral chemical, andhydrologic information based on field and laboratory studies are presented in order todistinguish mineral assemblages and textures associated with three paragenetic stages ofalteration at Teigarhorn: syn-eruptive, near surface alteration (Stage I), regional burialmetamorphism (Stage II), and local hydrothermal alteration associated with dike emplace-ment (Stage III). Each stage is characterized by distinct mineral parageneses andtextures, with Stage II resulting in formation of the regional zeolite zones described byWalker (1960b). Bulk porosity varied considerably during the metamorphic history ofthe lavas due to filling of pore spaces by secondary minerals and generation of secondaryporosity through brittle deformation. The results provide an example of how observa-tions of mineral paragenesis and porosity evolution can be integrated in a petrotectonicinterpretation of low-grade metamorphism in basaltic terraines.

GEOLOGICAL SETTING

Geology.—Teigarhorn is located near the eastern edge of the volcanic plateau ofpresent-day Iceland (fig. 1). Lavas erupted from fissures and central volcanoes duringmagmatism along a divergent plate boundary between the North American and Eur-asian plates (Kristjansson, Gudmundsson, and Haraldsson, 1995). The oldest exposedlavas are only about 13 Ma (Watkins and Walker, 1977). Upper crustal stratigraphy ineastern Iceland consists largely of subaerial basaltic lavas with minor intercalatedsediments, silicic lavas, lignite layers, and pyroclastic units (Robinson and others, 1982;Gustafsson, 1992; Walker, 1960b; Wood, 1977; Sæmundsson, 1979). Continued volca-nism and rifting led to pronounced tilting of the lavas that dominate crustal structurearound Teigarhorn (Palmason, 1973); dips increase from �2° to 6° southwest at the topof the lava pile to �6° to 9° southwest at sealevel (fig. 1; Walker, 1974).

Disrupting the monoclinal structure of the lava pile are mafic and silicic intrusivecenters and central volcanic complexes (fig. 1; Walker, 1963). Extrusive products of theextinct central volcanoes are interlayered with fissure-erupted lavas that make up most ofthe crust in eastern Iceland. Glacial exhumation has revealed an alignment of mafic diceswarms, central volcanoes, and intrusive centers that defines the geometry of paleo-riftzones. Teigarhorn lies within a prominent dike swarm (called the Alftafjordur dike

Philip S. Neuhoff and others—Porosity evolution and mineral468

swarm; Walker, 1963): the Alftafjordur volcanic center and the Austurhorn intrusionboth lie within this swarm (figs. 1 and 2; Walker, 1963; Blake, 1970). The compositionand mineralogy of these dikes are similar to the surrounding basalts discussed below(Walker, 1963). The dikes are subvertical, trend north-northeast (fig. 2), and vary inthickness from 0.25 to 3 m.

Lithology.—The volcanic stratigraphy exposed at Teigarhorn includes approx 150 mof subaerial mafic lava flows with minor intercalated pyroclastic units. The uppermost

Fig. 1. Generalized geological map of a portion of eastern Iceland (after Johannesson and Sæmundsson,1989) showing the distribution of major rock types, locations of intrusions and central volcanoes, and selectedlava orientations from Walker (1963, 1974). Dashed curve outlines the Alftafjordur mafic dike swarm wheredikes make up greater than 8 percent (by volume) of the crust (Walker, 1963).

paragenesis during low-grade metamorphism of basaltic lavas at Teigarhorn 469

Fig. 2. Map of Teigarhorn showing the distribution of individual mafic dikes (determined by fieldmapping and from aerial photographs) and the locations of samples and porosity measurements. Also shown isthe location of the Skessa tuff (Walker, 1963).

lavas belong to the Grænavatn Porphyritic Group (Walker, 1959), and the oldest unitsare the Skessa tuff (see below) and surrounding undesignated lavas. These units wereburied to approx 1500m at Teigarhorn (Walker, 1960b). Compositionally, the lavasexposed at Teigarhorn (fig. 2) are 3 to 25 m thick subaerial low-MgO tholeiites andferrobasalts (Flower and others, 1982). Most flows are fine-grained, containing abundantgroundmass plagioclase, olivine, pyroxene, and glass with minor olivine and pyroxenephenocrysts. Lavas with up to 30 modal percent plagioclase phenocrysts occur infre-quently. The flows have scoraceous, brecciated flow tops. Vesicles are abundant near thetop and bottom of the flows, but the flow centers are essentially non-vesicular andmassive. A 20 m thick, green, welded, silicic tuffaceous unit (the Skessa tuff of Walker,1959) is exposed near the southern portion of Teigarhorn (marked ‘‘Skessa tuff’’ on fig.2). The Skessa tuff is a regionally extensive unit that appears to be the product of aneruption from an unexposed volcanic center approx 25 km north of Teigarhorn (Walker,1962). Separating many flows are undulatory weathering surfaces that frequently areoverlain by bright red basaltic tephras that are 1 to 15 cm thick and constitute less than 1percent of the stratigraphy. The tephras are phreatomagmatic deposits produced byinteraction between basaltic magmas and surface or groundwaters at eruption sites 10’sto 100’s of km away from Teigarhorn (Heister, ms).

The volcanic stratigraphy in eastern Iceland has been subjected to regional zeolite-facies metamorphism; higher grades of metamorphism are exposed around intrusivecomplexes and central volcanoes (Walker, 1960b). Walker (1960b) described depth-controlled zeolite-facies mineral zones (fig. 3) that parallel the paleosurface of the lavapile and transgress the volcanic stratigraphy. This implies that the zones formed aftertilting of the lava pile and do not solely reflect variations in chemistry of the host lavas.The zeolite zones described by Walker (1960b) are underlain by zones characterized bystilbite � heulandite � mordenite, laumontite, and prehnite � epidote with increasingdepth (Mehegan, Robinson, and Delany, 1982). Coastal exposures around Berufjordur(fig. 1) have provided museum-grade zeolite specimens for over 200 yrs (Betz, 1981).Teigarhorn, in particular, has been an important source for zeolites used in studies ofcrystal chemistry (Galli and Rinaldi, 1974; Passaglia, 1975; Passaglia and others, 1978;Alberti, Pongiluppi, and Vezzalini, 1982), crystal structures (Slaughter and Kane, 1969),reaction kinetics (Ragnarsdottir, 1993), optical properties (Bauer and Hrıchova, 1962),and piezoelectricity (Bellanca, 1940).

METHODS

Mapping and sample collection.—Representative samples of altered and unaltered lavaswere collected throughout the coastal exposures at Teigarhorn in conjunction withmapping of dike distribution and outcrop porosity measurements (fig. 2). Estimates oftotal macroscopic porosity (herein defined as the percentage of a given volume of rockoccupied by pore space) prior to mineral precipitation were conducted in situ duringmapping following the methods of Manning and Bird (1991). In the present communica-tion, a distinction is made between the total porosity at a given point in time (�) and thecumulative porosity that reflects pore abundance (�*) but does not include the effects ofpore occlusion by mineral precipitation (Manning and Bird, 1991). The contribution ofprimary pores to �* (�*1) was estimated by placing a clear mylar sheet imprinted with a54 by 41.5 cm grid (node spacing 1.5 cm) directly on the outcrop as a basis for pointcounting macroscopic pores. Estimates of �*1 included vesicles and other primarystructures as well as open space between flow units that could not texturally be linked topost-eruption processes. Although this technique only identifies macroscopic pores onthe order of 1 mm or larger, the contribution of these pores to �*1 is large, anduncertainties are on the order of 1.5 percent or less for grids of over 1000 points(Manning and Bird, 1991; Chayes, 1956). The contribution of secondary pore spaces to

Philip S. Neuhoff and others 471

�* (�*2) was estimated by measuring cross-sectional areas of fractures/veins, breccias,and dissolution features associated with late-stage alteration within 3 to 6 m2 sections ofoutcrop. Although outcrop area measurements of pore space are scale dependent andnot necessarily representative of bulk properties (Manning and Bird, 1991), they areuseful for assessing the general magnitude and variability of �*2. Repeated area measure-ments indicate that uncertainties were �1 percent.

Analytical techniques.—Petrographic and hand-sample observations were conductedon 20 samples representative of lithologies and mineral parageneses at Teigarhorn.Optical properties of minerals in transmitted light were used along with mineral habits inboth hand specimen and thin section to augment field observations of mineral paragen-eses and phase relations. Estimates of the percentage of pore space occupied by variousalteration minerals were conducted by digital analysis of photomicrographs. Images ofdiscrete pores were digitized using the software package Scion Image (Scion Corpora-tion, Frederick, Maryland) to calculate the areas of various secondary phase/assemblages and residual porosity relative to the total area of the pore. Repeat measure-ments indicate that errors in area measurements were less than 1 percent.

Fig. 3. Depth distribution of regional zeolite facies mineral zones in eastern Iceland. Zones above sealevelare from mapping in olivine normative tholeiites around Berufjordur by Walker (1960b) and below sealevel arefrom studies of the Reydarfjordur drill core located �40 km north-northwest of Teigarhorn (Mehegan,Robinson, and Delaney, 1982). Note that the outcrops at Teigarhorn are approximately at sealevel. Stage IIalteration at Teigarhorn includes both the mesolite/scolecite and heulandite � stilbite mineral zones.


A battery of analytical techniques was used to confirm phase identities andcompositions. Powder X-ray Diffraction (XRD) was conducted on mineral separatesusing a Rigaku theta-theta diffractometer with CuK� operating at 35 kV and 15 mA.Mineral separates for XRD identification were extracted from individual vesicles with arotary drill. Mineral chemistry was determined by electron probe microanalysis (EPMA)on an automated JEOL 733A electron microprobe operated at 15 kV acceleratingpotential and 15 nA beam current. Calibration was conducted using natural geologicstandards. Beam width for analysis of hydrous minerals (i.e., that is, clays and zeolites)was varied between 10 and 30 µm depending on grain size. Raw counts were collectedfor 20 s and converted to oxide wt percents using the CITZAF correction procedure afteraccounting for unanalyzed oxygen following the methods of Tingle and others (1996).Raw counts for individual elements (including Na) were essentially constant with time,suggesting the elemental drift (loss) was negligible. Bulk rock compositions weredetermined on a Rigaku fully automated X-ray fluorescence (XRF) spectrometer. Fluidinclusion homogenization and freezing temperatures were determined on one sample ofcalcite (Iceland spar variety) using gas-flow heating/freezing stage adapted by Fluid, Inc.

RESULTS

Mineral ParagenesisAlteration at Teigarhorn occurred during three paragenetic stages. The textural and

mineralogical characteristics of these stages are summarized in table 1 and the sequenceof mineral precipitation during each stage is shown in figure 4. Representative composi-tions and molar formulas for clays, zeolites, and feldspars are given in tables 2 through 5.

Stage I.—Paragenetic stage I is the earliest alteration at Teigarhorn and is character-ized by linings of primary pore space by celadonite and silica minerals (table 1; figs. 4and 5B). All other mineral assemblages are either precipitated on top of or cross cut StageI alteration (fig. 5B). Stage I alteration is common but not present in all samples.Celadonite and silica minerals typically form layers along the walls of pore space with alayered or botryoidal habit. Occasionally Stage I silica forms layered geopetal structures

TABLE 1

Mineralogical and textural characteristics of paragenetic stages observed at Teigarhorn,eastern Iceland

ParageneticStage*

Mineral Assemblagesin Pore Space

MineralAssemblages

in Matrix Textures

Stage I �celadonite � silica minor celadonite pores: lining primary porespace; geopetal structures

matrix: minor celadonitereplacing groundmass

Stage II C/S2 � scolecite � chalcedony �quartz

C/S � albite �zeolites

pores: C/S pore linings

C/S � stilbite � heulandite �mordenite � epistilbite �qu artz � chalcedony

partial or complete primarypore fillings by zeolites

matrix: pseudomorphicreplacement of groundmassand phenocrysts

StageIII quartz � clinoptilolite � stilbite �heulandite � mordenit e �scolecite � laumontite � calcite

quartz � C/S pores: brecciation, fracturing,mineral precipitation

matrix: discoloration (SiO2metasomatism)

*Stage I was the earliest stage, Stage III the last. 2Mixed-layer chlorite smectite (see text).


on the bottom of vesicles that are oriented roughly parallel to the lava stratigraphy.Minor celadonite occurs in the volcanic groundmass.

Celadonite typically precedes silica in the crystallization sequence (fig. 5B). Arepresentative celadonite composition from Teigarhorn is given in table 2, and 6celadonite analyses are plotted in figure 6. Stage I celadonite is close to the endmembercompositions as defined by Buckley and others (1978), with only slightly lower potas-sium content and higher aluminum than the ideal endmember composition[K

2(Fe,Al)2(Mg,Fe)2Si8O20 (OH)4; table 2; fig. 6]. The slight deviation from ideal compo-

sition may be due to the presence of a small trioctahedral smectite component (either asintergrowths or interstratifications) in the celadonites, although smectite is not detectedin the XRD patterns. Celadonite is distinguished from other phyllosilicates by itsdistinctive bright green color in both hand specimen and thin section (figs. 5B and 5D),its layered habit, and by its potassium-rich composition (table 2; fig. 6). Silica mineralsconsist of quartz and/or chalcedony. It appears likely that quartz and chalcedonyreplaced preexisting Stage I opaline silica during burial, as opaline silica is the only StageI silica phase observed near the top of the lava pile above Teigarhorn (Neuhoff,unpublished data).

Stage II.—Paragenetic Stage II alteration is characterized by continued infilling ofprimary pore spaces and partial replacement of groundmass and phenocryst minerals.As with Stage I, the distribution at the outcrop scale of Stage II alteration is largely afunction of the distribution of primary porosity in the lavas (fig. 5A). Stage II mineralassemblages and paragenetic sequences are summarized in table 1 and figure 4. Mixedlayer chlorite/smectite (C/S) is the first phase to precipitate in primary pore space duringStage II (fig. 5B, C). The term ‘‘mixed layer chlorite/smectite’’ and the abbreviation C/Sare used in the present communication without specific inference of ordering amonglayers (Reynolds, 1988). Brown to brownish green radiating bundles of C/S for linings ofprimary pore space. The bundles nucleated on pore walls or on Stage I minerals wherepresent and developed toward the center of the pores. As shown in figure 6, Stage II C/S

Fig. 4. Temporal sequence of secondary minerals developed in lavas at Teigarhorn. Events toward theright of the diagram occurred after events to the left.


compositions (table 2) project near the binary compositional join between ferromagne-sian trioctahedral smectite (saponite-ferrosaponite; Guven, 1988) to ferromagnesianchlorite (clinochlore-chamosite; Bailey, 1988). The percentage of chlorite (or smectite)interlayers was estimated from electron microprobe analysis of the total molar Si � Al �Mg � Fe content of the mineral by linear interpolation between smectite and chloriteendmembers (Schiffman and Fridleifsson, 1991). This interpolation requires calculationof molar formulas based on 28 oxygen charge equivalents (O � 2OH), for which themolar Si � Al � Mg � Fe contents of endmember chlorite and trioctahedral smectite are20 and 17.82, respectively. By this analysis, Stage II clays contain between 12 and 85 molpercent chlorite. The most prominent interlayer cations in C/S clays are Ca�2 and Na�

(table 2). Potassium typically accounts for less than 5 mol percent of the total interlayercations but in rare cases constitutes more than 25 mol percent with one sample as high as60 mol percent. Mixed-layer chlorite-smectite with high K� contents always occur inclose proximity to celadonite, and textural evidence suggests that local replacement ofStage I celadonite by Stage II C/S may have occurred.

The compositions of C/S clays become progressively enriched in chlorite duringStage II alteration. Figure 7 shows the percentage of chlorite in a rim of Stage II C/S as afunction of distance from the vesicle wall. In general, C/S near the wall of the vesicleprecipitated earlier than that closer to the center of the vesicle. It can be seen in figure 7that the chlorite content of the clay generally increases from �20 percent at the vesiclewall to about 80 percent near the center of the vesicle, indicating that the chlorite contentof Stage II C/S increased with time.

The last phases to crystallize during Stage II are minor quartz and chalcedony alongwith zeolites (table 1 and fig. 4) that serve as index minerals for the regional zeolite zonesshown in figure 3. Zeolites partially or completely fill the pore space remaining afterprecipitation of C/S (fig. 5B). As indicated in table 1, two Stage II zeolite assemblagesoccur at Teigarhorn; the mesolite � scolecite zone described by Walker (1960b; fig. 3) isunderlain by a zone characterized by heulandite � stilbite with occasional mordeniteand epistilbite. The lower zone is similar to zones characterized by heulandite and/orstilbite identified in Icelandic geothermal systems (Kristmannsdottir and Tomasson,1978), East Greenland (Neuhoff and others, 1997), the Faeroe Islands (Jørgensen, 1984),the North Shore Volcanic Group (Minnesota, USA; Schmidt, 1990), and the Reydarfjor-dur drillhole (Mehegan, Robinson, and Delaney, 1982).

Mesolite � scolecite zone alteration at Teigarhorn is characterized by radiatingbundles of near stoichiometric scolecite fibers (table 3) that fill nearly all pore spacesremaining after C/S precipitation (fig. 5D). Scolecite is white in hand specimen and oftenhas a brownish hue in plane polarized light. Growth direction is very distinct; thebundles widen as a function of distance from the apex where nucleation first occurred atthe edge of Stage II C/S pore linings.

Typical vesicle fillings in the heulandite � stilbite zone consist of intergrowths ofstilbite, heulandite, and mordenite with occasional epistilbite (fig. 5B, C). Frequentlyonly one or two of these minerals are present in a given pore (usually stilbite orheulandite). Mordenite is fibrous, occurring alone or as discrete fibers within heulanditeor stilbite crystals (fig. 5C). Stilbite, heulandite, and epistilbite occur as euhedral blockyto lath shaped crystals. No evidence was found for heulandite � stilbite zone assem-blages overprinting (replacing) mesolite � scolecite zone minerals. Representativecompositions of Stage II stilbite, heulandite, and epistilbite are given in table 3 (Stage IImordenite is too fine-grained for reliable analysis). Epistilbite compositions are slightlyricher in Al�3 and Na� than the ideal formula (CaAl2Si6O16 · 6H2O). The ratio of Si to Alin Stage II heulandite and stilbite ranges from �2.9 to �3.4 and �2.9 to �3.5,respectively. Exchangeable cation sites in Stage II heulandite (fig. 8) and stilbite (fig. 9)



Fig. 5. Stage I and II alteration at Teigarhorn. (A) Outcrop view showing the flow scale distributionof primary pore spaces that host Stage I and II alteration. Two lava flows are visible separated by an undulatingsurface. The top of the lower flow is highly vesicular and contains abundant Stage I celadonite rims and Stage IIscolecite visible as white pore fillings. The massive flow center of the upper flow contains only isolated,large vesicles filled with Stage II scolecite. (B) Stage I celadonite and silica rims preceding Stage II mixed-layer chlorite/smectite (C/S) and stilbite in sample ISMP-2. Photomicrograph taken with polars partiallycrossed. Analyses in columns 2 and 7 of table 2 were conducted on celadonite and C/S rims shown in photo.

(C) Stage II heulandite � stilbite zone assemblages (sample 94-53). Photomicrograph taken with polarspartially crossed. Vesicle in rimmed with celadonite (Stage I) and Stage II C/S. Filling the pore are lath-shapedcrystals of stilbite (birefringement phase) and blocky crystals of heulandite. The stilbite and heulandite crystalscontain abundant mordenite fibers. Blue areas are residual porosity (color due to impregnated epoxy). (D)Stage II mesolite � scolecite zone assemblages (sample 94-77). Photomicrographs taken with polars partiallycrossed. Vesicles are rimmed with C/S and filled with scolecite. The large plagioclase phenocrysts have beenreplaced by scolecite.


are dominantly occupied by Ca�2, although in some samples Na� and K� constitute upto 30 percent of the exchangeable ions.

Stage II minerals in the matrix of the lavas occur predominantly as pseudomorphicreplacements of primary phenocryst and groundmass phases (table 1). Groundmassglass, pyroxene, and olivine as well as pyroxene and olivine phenocrysts are partiallyreplaced by C/S with compositions similar to that lining primary pore space. Ground-mass and phenocryst plagioclase is pseudomorphically replaced by scolecite and/orcompletely albitized (table 4).

Stage III.—The latest alteration at Teigarhorn (Stage III) is developed in secondarypore space (fractures, breccias, et cetera) and primary pores intersected by this porosity(figs. 10, 11). Stage I and II minerals filling vesicles and other primary pore spaces arecross cut by Stage III porosity. Stage III breccias commonly entrain Stage I and IIminerals (fig. 11C). This is frequently visible in outcrop as rafts of vesicle walls coatedwith celadonite and Stage II heulandite within Stage III breccias. Development offracture porosity during Stage III is associated with intrusion of mafic dikes that cross cutthe lavas. The near-vertical dikes make up more than 8 percent of the volume of rocks atTeigarhorn (fig. 2) and vary in thickness between 0.5 and 3 m. Spacing between dikes isuneven, ranging from less than 1 to over 30 m. Lava flows within a few meters of thedikes are extensively fractured, but the fractures generally do not contain secondary

Fig. 6. Chemistry of phyllosilicates from Teigarhorn (normalized to 28 O charge equivalents) as a functionof total mols of charge associated with exchangeable ions (2Ca � Na � K) and total molar Si � Al � Mg � Fecontent. Stage I phyllosilicates are shown as solid diamonds, Stage II phyllosilicates as open squares, and StageIII phyllosilicates as solid circles. Ideal, endmember celadonite, chlorite, and trioctahedral smectite are plottedas large black ovals. Note restricted composition ranges for Stage I and III phyllosilicates and higher smectitecontents of Stage II phyllosilicates.


TA

BL

E2

Rep

rese

ntat

ivep

hyllo

silic

atea

naly

sisfr

omTe

igar

horn

,eas

tern

Icel

and

dete

rmin

edby

EPM

A

1 oxy

gen

char

geeq

uiva

lent

spe

rfo

rmul

aun

it2 a

llir

onas

sum

edto

beFe

�2


minerals. At distances greater than 3 m from some dikes, brittle deformation featureshost Stage III mineral assemblages. This alteration is only locally developed.

Stage III mineral assemblages fill 0.2 to 2 cm thick veins oriented parallel to thedikes and irregular fractures and breccias up to 0.8 m across. Veins and fractures can betraced for up to 10 m. The margins of veins are occasionally lined with 5 micron equantcrystals of clinoptilolite (compositions given in fig. 8 and table 5) and usually marked byone or more layers of chalcedony. Filling the veins is quartz that is an- to subhedral nearthe margin and euhedral in the center with the c axis oriented normal to the vein walls.Clusters of euhedral quartz crystals (individual crystal size two to four mm) frequentlyform stalactiform deposits hanging vertically from the roofs of large cavities in thebreccias. Generally chlorite-rich C/S occurs intergrown with quartz and as discretefracture fillings (table 2). The last assemblages to fill the veins and breccias are euhedralstilbite, heulandite, or stilbite � heulandite � mordenite (fig. 11A, B) with localdevelopment of later scolecite, laumontite, and calcite in large, open cavities (fig. 11B).Stage III zeolites frequently occur as vein fillings and as large (cm-scale) euhedral crystalsfor which Teigarhorn is famous. Representative compositions of Stage III zeolites aregiven in table 5. The compositions of heulandite, stilbite, and scolecite formed duringStage III are similar to those formed during Stage II (table 5; figs. 8 and 9).

Fluid inclusion heating/freezing experiments were conducted on a 3 cm crystal ofIceland spar type calcite that formed late in Stage III along with scolecite and laumontite.Two-phase inclusions (liquid with vapor bubbles constituting �5 percent of the inclusion

Fig. 7. Temporal changes in the compositions of Stage II mixed-layer chlorite-smectite (C/S) depicted as afunction of percent chlorite (calculated using method of Schiffman and Fridleifsson, 1991). Analyses are from a300 µm linear traverse through a Stage II clay rim of a vesicle in sample 94-50. The compositions of C/S trendfrom early smectitic clays (near the vesicle wall) to relatively late chloritic clays (towards the center of thevesicle).


volume) occur throughout the crystal. Most inclusions are up to 0.25 mm in diameterwith ovoid, elongate, or triangular morphologies. Primary inclusions homogenized at115° � 8°C and frozen at �0.1° to 0°C.

Stage III fractures and breccias are frequently surrounded by aureoles (1-40 cmwide) of extensive SiO2 metasomatism (fig. 10). These zones are lighter in color than thesurrounding lavas, and in extreme cases are pale yellow. The lava is extensively replacedby fine-grained aggregates of quartz with about 5 modal percent C/S (fig. 11C and D).Stage III C/S (table 2, fig. 6) is more chloritic than in Stage II, ranging from 55 to 90percent chlorite. Where this alteration is most intense the original texture of the lava iscompletely obliterated. The replaced lava is highly porous with a spongy texture (fig.11D) and has approx 20 percent secondary porosity.

Bulk Rock ChemistryBulk rock compositions of samples from two successive lava flows (and an interven-

ing mafic volcanoclastic unit) from Teigarhorn (samples 94-86, 94-87, 94-90, 94-91,94-96, and 94-101; fig. 2) are given in table 6. Samples 94-86 and 94-101 are from theoverlying (younger) flow, 94-87 is the mafic volcanoclastic horizon, and 94-90, 94-91,and 94-96 are from the underlying (older) flow (see footnote to table 6 for sampledescriptions). The least altered samples of both flows (94-86 and 94-90) are similar incomposition and belong to the low MgO tholeiite lavas described by Flower and others(1982). The mafic volcanoclastic unit is compositionally similar to the unaltered lavasexcept for higher Fe, Si, K, Rb, Cu, Cr, and Ni and lower Ca, Na, P, Sr, and V contents inthe oxide-rich, more oxidized volcanoclastic horizon.

The effects of alteration on bulk rock composition can be assessed by comparingthese relatively unaltered samples with examples of Stages I and II (94-91 and 94-96) andStage III (94-101) alteration from the same flows. This is illustrated by the isocon plots

TABLE 3

Representative compositions of Stage II zeolites from vesicles determined by EPMA


(Grant, 1982) in figure 12, which compare the compositions of sample 94-96 (whichexhibits Stage I and II alteration) and sample 94-101 (extreme Stage III SiO2 metasoma-tism) to those of the relatively unaltered samples 94-90 and 94-86, respectively, of thesame flows. The linear trends (labeled ‘‘constant mass isocon’’ in fig. 12) represent noelemental loss or gain during alteration. It can be seen in figure 12A that Stage I and IIalteration together lead to only minor changes in bulk rock composition, as illustrated bythe fact that the major element data plot along the isocon. With the exception of Rb andLa in sample 94-96, the other trace element data listed in table 6 have been omitted forclarity, but they all plot along the constant mass trend as well. The most significant bulkchemical signature of Stages I and II is an increase in La, Rb, and to a lesser extent K2Oconcentration; Rb and K2O were likely concentrated due to precipitation of celadonitein 94-96. The fact that most of the data in figure 12A plot along the constant mass isoconreflects the isochemical nature of the hydration reactions that typify Stage I and IIalteration. In contrast, the extreme metasomatism associated with Stage III is clearlyreflected in sample 94-101. Quartz along with minor stilbite and chlorite have entirelyreplaced the lava (94-86) resulting in a bulk rock composition that is �92 percent SiO2.As can be seen in figure 12B, this sample is clearly enriched in SiO2 and depleted in allother analyzed elements relative to the unaltered lavas.

PorosityLava flows at Teigarhorn exhibit values of �*1 (volume percent of primary pores) up

to 22 percent with subsequent processes during Stage III leading to values of �*2 of up to11 percent (table 7). The range of values of �*1 given in table 7 is typical of basaltic lavas

Fig. 8. Molar proportions of Ca, Na, and K in heulandite and clinoptilolite from Teigarhorn.


reported by Christensen and Wilkens (1982) from the Reydarfjordur drillhole in easternIceland. Variations in �*1 (table 7) reflect differences in pore density both between flowsand within various portions of a given flow. The estimates of �*2 (volume percent ofsecondary pores) are similar to those obtained by Manning and Bird (1991) for fracturedbasalts near the margin of the Skaergaard intrusion in East Greenland.

Percentages of �*1 occupied by Stage I celadonite and silica and Stage II clays andzeolites along with the percent of �*1 remaining after precipitation of Stage I and IIminerals are given in table 8. Inspection of the table shows that the extent of porosityreduction during surficial alteration and burial metamorphism is heterogeneous betweensamples and is a strong function of paragenetic stage. Precipitation of Stage I minerals,where developed, leads to a 10 to 30 percent reduction in �*1. Chlorite/smectiteinterlayer clays are more common than Stage I minerals and fill 0 to 80 percent of �*1.Zeolites, the last minerals to precipitate during Stage II, fill an additional 0 to 90 percentof �*1. The percent of �*1 remaining after Stage II varies between �0 and 35 percent; thatis, � after Stage II ranges from about 0 percent to nearly 8 percent.

DISCUSSION

Time-space development of mineral parageneses.—Mineral parageneses developed atTeigarhorn present a time-integrated view of the petrogenesis of zeolite-facies basalticaquifers in eastern Iceland. In this context, the temporal development of mineralassemblages can be directly related to the geologic evolution of the crust in easternIceland. This is summarized in figure 13, which depicts the history of the lavas at

Fig. 9. Molar proportions of Ca, Na, and K in stilbite from Teigarhorn.


Teigarhorn. The horizontal axis in figure 13 represents time after eruption and thevertical axis is crustal depth. Also illustrated in figure 13 are petrotectonic events (lavaeruption, burial during crustal accretion, dike intrusion, glacial exhumation) and mineralparageneses experienced by the lavas.

Celadonite and silica rims (Stage I) formed at shallow depths (fig. 13). Stage Igeopetal structures at Teigarhorn have orientations similar to the lava flows, and in otherlocalities exhibit progressively more shallow dips between silica bands due to progres-sive tilting (Walker, 1974). This relationship suggests that the geopetal structures formedprior to tilting. Near surface genesis for Stage I minerals is consistent with the lowtemperatures (� �50°C) at which celadonite forms in natural systems (Odin, 1988) andthe supersaturation of Icelandic surface waters with respect to chalcedony (Gıslason andEugster, 1987b; Gıslason, Arnorsson, and Armannsson, 1996). Silica and potassium areamong the most mobile elements in near-surface environments in Iceland (Gıslason,Arnorsson, and Armannsson, 1996), providing components necessary for formation ofceladonite and opal.

TABLE 4

Representative compositions of altered and fresh plagioclasedetermined by EPMA


Burial metamorphism resulted in formation of Stage II mineral assemblages.Mixed-layer chlorite/smectite clays most likely formed during burial (fig. 13). Chemicalcomponents necessary for C/S precipitation were derived from hydrolysis of olivine andbasaltic glass. We propose that the increase in chlorite content of C/S pore linings withcrystallization time (fig. 7) is a consequence of progressive growth of C/S as temperatureincreased during burial of the lavas. The trend toward endmember chlorite compositionshown in figure 7 is analogous to prograde zones characterized by smectite, C/S, andchlorite observed with increasing temperature in active geothermal systems in Iceland(Kristmannsdottir, 1979; Schiffman and Fridleifsson, 1991). The early, more smectite-rich C/S formed at relatively low temperatures (shallow depths) and as burial proceededC/S became progressively more chlorite-rich as temperature increased during burial.

Precipitation of Stage II zeolites occurred after burial of the lavas. Regional mineralisograds in eastern Iceland characterized by Stage II zeolites, such as the boundarybetween the mesolite � scolecite zone and the heulandite � stilbite zone, transgress(crosscut) the tilted volcanic stratigraphy in eastern Iceland (Walker, 1960b). The zeoliteisograds thus postdate tilting of the lava pile, as observed in the Early Tertiary plateaubasalts of East Greenland (Neuhoff and others, 1997). In East Greenland, the sequence ofregional zeolite mineral zones with depth did not develop successively as the lavas wereburied. Instead, it appears that regional zeolite formation occurred after the main stage ofEarly Tertiary volcanism in East Greenland ceased. The notable lack of scoleciteprecursors to heulandite � stilbite zone mineral assemblages at Teigarhorn suggest asimilar model for the timing of zeolite zone development in eastern Iceland. Zeolites didnot precipitate during burial of the lavas. The pervasive albitization observed at

Fig. 10. Stage III fractures filled with quartz (outcrop view). Pencil for scale. The dashed white lines areboundaries of metasomatic zones containing quartz � minor C/S that replace the basalts (see fig. 11C and D).The aureoles are visible as areas of yellow alteration marked by the symbol ‘‘Si.’’ Also visible are vesiclescontaining Stage II scolecite.


Fig. 11. Texture and mineralogy of Stage III alteration at Teigarhorn. (A) Cavity in Stage III breccia linedwith euhedral stilbite. Pencil for scale. (B) Stage III breccia lined with heulandite and later stilbite and scolecite.Primary vesicles in lava around cavity are lined with green celadonite. Pencil for scale. (C) Timing and texturalrelationship between Stages I and III (sample 94-94). Photo taken with polars slightly crossed. The wall of aprimary pore space is lined with celadonite; Stage III stilbite fills a vein along the wall of the pore. A patchyaggregate of quartz (with minor C/S) is replacing the lava on the other side of the vein. The ‘‘pie slice’’—shapedceladonite in the center of the photo has been dislodged from the wall of the pore space during Stage IIIbrecciation. Note the vuggy texture of the quartz aggregate replacing the lava (blue areas are porosity filled withimpregnated epoxy). (D) Hydrothermally altered lava (sample 94-97) that has undergone extensive SiO2metasomatism. Photo taken with polars slightly crossed. The matrix of the lava is pervasively replaced byquartz and C/S. Stilbite filled cavities appear to represent both partially replaced Stage II vesicle fillings andlate Stage III veins. Abundant pore space is present between quartz grains (blue areas).

Teigarhorn is not present at elevations �100m in lavas exposed in the mountains aroundBerufjordur (Neuhoff, unpublished data). This suggests that albitization occurred concur-rently with Stage II zeolite formation. The paucity of absolute age dates in easternIceland do not permit concise constraints on the timing and duration of zeolite

Fig. 11 (continued)

Philip S. Neuhoff and others 487

TA

BL

E5

Rep

rese

ntat

ivec

ompo

sitio

nsof

Stag

eIII

frac

ture

-filli

ngze

olite

sdet

erm

ined

byE

PMA

.


precipitation during Stage II; however, it seems likely that it did not last more than a fewMa, as observed in East Greenland (Neuhoff and others, 1997).

Pore structures filled with Stage III alteration cross-cut Stage II mineralization. StageIII is spatially and, presumably, temporally linked to dike intrusion (fig. 13). The spatialassociation of Stage III with late dikes indicates that this last paragenetic stage is a local,and not a regional, phenomenon. The close association with dike intrusion indicates thatStage III occurred while the region around Teigarhorn was still undergoing intrusive

TABLE 6

Bulk rock compositions from Teigarhorn, eastern Iceland

‡Sample descriptions:94-86: compact lava with minor Stage III heulandite veins94-87: bright red mafic volcanoclastic unit marking boundary between two lava flows94-90: compact, massive lava with minor small vesicles filled with smectite94-91: compact basalt with minor vesicles containing Stage II clay minerals94-96: vesicular lava with abundant Stage I celadonite, Stage II heulandite and smectite94-101: yellow, SiO2-metasomatized lava from aureole around Stage III breccia*Weight percent of oxide in anhydrous (calcined) sample. †Total Fe reported as Fe203. §Determined by loss onignition at 110°C (H2O�) and 900°C (H2O�). †Parts per million of element in untreated sample. B/D: belowdetection limit.


Fig. 12. Isocon plots (Grant, 1982) depicting the variation in bulk rock chemistry accompanying Stages Iand II (A) and Stage III (B) alteration at Teigarhorn. The lines labeled ‘‘constant mass isocon’’ depict thetheoretical distribution of data assuming that the mass and compositions of the lavas do not change duringalteration. The SiO2, Rb, and La analyses have been multiplied (by factors of 0.25, 10000, and 10000,respectively) to facilitate comparison with data for other major oxide components. Data used to construct thisplot are listed in table 6.


activity related to rifting. Increased heat flow and permeability due to fracturing duringdike emplacement initiated hydrothermal fluid circulation that led to the extensivewall-rock replacement and fracture filling characteristic of Stage III mineralization. Noage dating has been conducted on the dikes at Teigarhorn that would permit constraintson the absolute timing of Stage III alteration.

Thermobarometric conditions during Stages II and III.—Barometric conditions duringparagenetic Stages II and III were complex functions of the burial and hydrogeologic

TABLE 7

Outcrop porosity measurements

1Station locations given in figure 2. Descriptions: (1) flow top; (2) flow center with abundant Stage III fracturesdeveloped �10 m from dike; (3) flow top; (4) flow center; (5) flow center; (6) flow top; (7) flow center withabundant Stage III fractures developed �15 m from dike; (8) flow bottom with abundant Stage III fracturesdeveloped about 5 m from dike.

TABLE 8

Percent of pore space occupied by Stage I and II minerals based on thin section area measurement.All samples are from flow tops

aNumber of individual pores measured. bRims of paragenetic Stage I celadonite or silica on vesicle walls.cIncludes all paragenetic Stage II trioctahedral smectite and chlorite interlayer clays. dStage II zeoliteassemblages, dominantly mesolite � scolecite and heulandite � stilbite � mordenite. ePrimary pore space (�*1)remaining after paragenetic Stage II.


history. Pressures during precipitation of Stage II C/S increased from a minimum of 1bar (atmospheric pressure) to pressures between �170 and �500 bar, corresponding,respectively, to hydrostatic and lithostatic pressures at the maximum burial depth of1500 m for the lavas at Teigarhorn (Walker, 1960b). Occlusion of macroscopic porosityby Stage II clay and zeolite minerals (table 8) likely resulted in near-lithostatic pressures.Pressures during Stage III were probably closer to hydrostatic conditions due to thehigher permeability of continuous fractures (see below).

Although the systematic changes in Stage II C/S compositions appear to record thethermal history of the lavas (see above), the occurrence of this phase along with scolecite,stilbite, heulandite, and other Stage II zeolites is in contrast with the distribution of theseminerals in active Icelandic geothermal systems. Smectite is present at low temperaturesin Icelandic geothermal systems (below 200°C), with C/S occurring between 200° and220°C (Kristmannsdottir, 1979) and in some fields up to 270°C (Schiffman and Fridleifs-son, 1991). At higher temperatures chlorite is present. Thus, C/S in Icelandic geothermalsystems occurs at temperatures �100°C higher than the boundary between the mesolite �scolecite zone and stilbite zones (90° � 10 °C; Kristmannsdottir and Tomasson, 1978).

In regionally metamorphosed basalts, the boundary between zones characterizedby smectite and C/S is generally observed under zeolite-facies conditions (Schmidt andRobinson, 1997; Robinson, Bevins, and Rowbotham, 1993; Bevins, Robinson, andRowbotham, 1991). For example, in the North Shore volcanic group of Minnesota,smectite coexists with heulandite and stilbite; C/S is associated with higher gradelaumontite-albite assemblages (Schmidt and Robinson, 1997). Based on the associationof C/S with scolecite and heulandite � stilbite at Teigarhorn, C/S formed at even lower

Fig. 13. Spatial and temporal development of pore-filling mineral assemblages at Teigarhorn. The verticalaxis depicts depth below land surface at the time of each event depicted in the figure. Time elapsed aftereruption increases to the right. No scale is implied on the horizontal axis. The parallel curves on the figure areboundaries of the lavas exposed at Teigarhorn in space and time. Note that the timing of dike emplacement ismeant to infer only the timing of dikes associated with Stage III alteration at Teigarhorn.


grades during Stage II in eastern Iceland. The relatively high chlorite contents of C/Sassociated with the zeolite assemblages at Teigarhorn may be a consequence of albitiza-tion of primary plagioclases, which promotes chlorite formation through release ofaluminum (Shau and Peacor, 1992; Schmidt and Robinson, 1997). As noted by Schiff-man and Fridleifsson (1991), the temperatures at which C/S occurs in geothermalsystems are often higher than in regional diagenetic systems. In contrast, temperatureestimates based on zeolite assemblages and fluid inclusion thermometry are similar (seebelow). This suggests that estimation of temperatures from observations of zeoliteassemblages in geothermal systems may be more reliable than similar estimates based onphyllosilicate distribution.

Assuming a temperature of formation for Stage II zeolite assemblages of 90° � 10°C(see above) leads to an estimated thermal gradient of approx 53° � 7°C/km (given asurface temperature of 10°C). A thermal conductivity for water-saturated basalt of 1.5 to2.0 W/(mK) (Oxburgh and Agrell, 1982) gives an estimated heat flow during Stage IIzeolite precipitation of 1.9 to 3.2 heat flow units (HFU � 10�6 cal · cm�2 · s�1). This valueis consistent with conditions on the flank of the present-day active volcanic zone inIceland (Flovenz and Sæmundsson, 1993).

Mineral assemblages and fluid inclusion measurements indicate that temperaturesduring Stage III were slightly higher than in Stage II. Empirically-based chloritegeothermometry (Cathelineau, 1988; Cathelineau and Nieva, 1985) suggests that therelatively early, chlorite-rich Stage III C/S formed at temperatures between 140° and180°C. It should be noted that temperatures estimated using chlorite geothermometrygenerally disagree with measured temperatures in Icelandic geothermal systems (Schiff-man and Fridleifsson, 1991). Late Stage III zeolite assemblages are broadly similar tothose in Stage II but may have formed at slightly higher temperatures based on thepresence of laumontite. Laumontite becomes prominent at temperatures above 120°C inIcelandic geothermal systems (Kristmannsdottir and Tomasson, 1978), consistent withpressure corrected homogenization temperatures for late-Stage III calcite (125° � 8°Cassuming hydrostatic pressures). The apparent decrease in temperature during Stage IIImay reflect dissipation of heat flow associated with crystallization of the dikes or may bethe result of errors associated with the chlorite geothermometer noted above.

Chemical fluxes during alteration.—The widespread precipitation of phyllosilicates,silica minerals, and zeolites in pore spaces during Stages I and II indicates that Si, Al, Ca,Na, K, Fe, and Mg were mobile over scales of a millimeter or more. This is furtherillustrated by the textural preservation of pore walls and by successive lining of porespaces during mineral paragenesis, which require that chemical components necessaryfor precipitation of secondary minerals were transported into the pores. All componentsnecessary to form the mineral assemblages described above were readily available in theprimary phases in the lavas (glass, olivine, plagioclase, pyroxene) with the exception ofH2O, which was provided by the metamorphic fluids. The similarity between the bulkrock compositions of lavas altered during Stages I and II and those of unaltered lavas(table 6; fig. 12A) indicates that many elements were relatively immobile (that is,conserved) on the scale of individual lava flows during regional alteration. An exceptionis potassium, which is concentrated during Stage I in celadonite (table 6). As notedabove, potassium is very mobile during near-surface alteration of basaltic lavas (Gısla-son, Arnorsson, and Armannsson, 1996). The concentration of Rb, K2O, and Laapparent in figure 12A and table 6 is consistent with the observations of Wood (1976)that these elements were mobilized during zeolite-facies metamorphism.

In contrast, high fluid fluxes through some Stage III (see below) fractures resulted inwidespread leaching at the outcrop scale of all elements except silicon (table 6; fig. 12B).Extensive replacement of the lavas by quartz (plus minor chlorite-rich C/S) is suggestiveof interaction with acidic fluids. The low pH of these fluids may reflect the presence of


relatively large amounts of dissolved magmatic gases. The solubilities of primary phasesin the lavas as well as C/S and zeolites are all enhanced under very acidic conditions. Incontrast, quartz solubility is insensitive to pH under acidic conditions, and it formedearly during Stage III from SiO2 released during hydrolysis/dissolution of primaryphases. The large model abundance (�90 percent) of quartz in Stage III aureolessuggests that some SiO2 may have been added during alteration. Aside from minor C/Sprecipitated in the aureoles and later in veins, no mineralogical sinks for iron andmagnesium formed during Stage III, suggesting that these elements were transportedover distances of meters (the scale of outcrops at Teigarhorn) away from their sources.Late precipitation of zeolites and calcite in Stage III veins most likely occurred as theaqueous solutions became more alkaline as protons were consumed by hydrolysis.

Both Stages II and III exhibit a pronounced temporal change during paragenesisfrom phyllosilicates � quartz to assemblages dominated by Ca-zeolites. In both cases,this paragenetic trend was probably coincident with changes in temperature andpressure. However, the abrupt shift in the mineral composition from C/S to zeolites mustreflect the chemical evolution of the aqueous electrolyte solutions from which theseminerals precipitated, as schematically illustrated in the phase diagram in figure 14. Thearrows A and B in figure 14 represent two possible trends in fluid composition that wouldresult in mineral paragenesis of Ca-zeolites after C/S (figs. 4, 5B and C). The decrease inlog [(aMg

�2/(aH�)2] suggested by paragenetic trends A and B would result from decreasingthe activity of Mg�2 in the pore fluid as a consequence of precipitation of C/S(represented by clinochlore in the phase diagram). Mineralogical sources of Mg�2

(olivine, basaltic glass) in basaltic lavas are generally less stable than plagioclase underlow-grade metamorphic conditions (Gıslason and Eugster, 1987a; Gıslason and Arnors-son, 1993) and are often hydrolyzed before plagioclase. Albitization of plagioclase waslikely contemporaneous with precipitation of zeolites in pores during Stage II (seeabove), adding Ca�2 to aqueous solution and increasing log ([aCa�2)/(aH�)2] as suggestedby trend A in figure 14. Instability of zeolites relative to C/S during early Stage IIalteration provides an explanation for the apparent lack of zeolite formation duringburial of the lavas inferred from our petrographic observations that heulandite � stilbitezone assemblages do not replace mesolite/scolecite zone assemblages. These observa-tions, combined with the phase relations depicted in figure 14, suggest that the lavas wereburied before olivine and basaltic glass were exhausted, resulting in fluid compositionsduring burial in which C/S was stable relative to zeolites.

The general paragenetic trend from C/S to zeolites, as well as fluid compositiontrends A and B in figure 14, is consistent with the results of experimental simulations ofbasalt-meteoric water interaction at low temperatures. Ghiara and others (1993) con-ducted experiments on closed-system dissolution of basaltic glass in deionized water at200°C that resulted in a mineral paragenesis of early smectite � phillipsite followed byanalcime. Total Mg concentrations in solution during these experiments initially roserapidly, and then declined to a steady state level as Mg-rich smectite precipitated.Calcium concentrations also rose rapidly during the initial stages of these experiments,and continued to increase until the last stages of the experiments. Thus, log [(aCa�2)/(aMg�2)] increased with time in response to precipitation of Mg-rich smectite and thefluids evolved toward the stability fields of zeolites as suggested in figure 14. Similarexperiments by Gıslason and Eugster (1987a) on dissolution of basaltic glass andcrystalline basalt at 25°, 45°, and 65°C resulted in nearly identical solution compositiontrends. Subsequent analysis of the solid phases produced in these experiments (Gıslason,Veblen, and Livi, 1993) identified several magnesium-bearing phases (kerolite, chryso-tile, talc, and smectite) along with quartz, calcite, kaolinite, and amorphous precipitates.Early formation of magnesian minerals in these experiments (including smectite) con-


sumed Mg, causing the value of log [(aCa�2)/(a Mg

�2)] to increase as inferred above fromnatural parageneses.

Fluid compositions in local equilibrium with Stage II mineral assemblages werepredicted using Version 8 of the computer code EQ3NR (Wolery, 1991) assumingheterogeneous equilibrium at 100°C, 500 bar between clinochlore, stilbite, quartz, andalbite. Additional constraints were provided by fixing the concentration of Cl� to 0.001(representative concentration in meteoric-water dominated geothermal systems in Ice-

Fig. 14. Stabilities of minerals in the system MgO-CaO-Na2O3-Al2O3-SiO 2-H2O-HCl at 100°C, 500 barin equilibrium with quartz and albite as functions of log [(aCa�2)/(aH�)2] and log [(aMg�2)/(aH�)2] in an aqueoussolution with aH2O � 1. Diagram calculated using thermodynamic properties and algorithms in the computercode SUPCRT92 of Johnson, Oelkers, and Helgeson (1992), with addition of data for stilbite(CaAl2Si7O18 · 7H2O; Gf

o (cal/mol) � �2381234; Hfo (cal/mol) � �25588306; So (cal/molK) � 191.41; Vo

(cm3/mol) � 332.61; Cpo (cal/molK) � 154.12 � 0.168 T(K) � 10860000 T(K)�2; Neuhoff, unpublished data).

The diagram was constructed for conditions approximating the maximum burial temperature and lithostaticpressure at Teigarhorn (see text). The right-hand axis gives values of log fco2

in equilibrium with calcite. Phaseboundaries between Al-bearing minerals are shown as solid lines. Saturation surfaces for stoichiometricdiposide, talc, wollastonite, and tremolite are depicted as dotted lines. Two sets of phase boundaries involvingclinochlore are depicted in figure 12: one for endmember clinochlore (aclinochlore � 1; solid lines) and onecorresponding to aclinochlore � 0.0068 (dash-dot-dot lines), the activity of clinochlore in sample 94-94 (table 2)assuming ideal mixing (Holland, Baker, and Powell, 1998). Prehnite is not depicted on the diagram as thethermodynamic properties reported for this phase in Johnson, Oeklers, and Helgeson (1992) imply that it isstable relative to all Ca-zeolites, conflicting with experimental and field evidence indicating that Ca-zeolites arestable relative to prehnite at this temperature and pressure. Trends A and B are discussed in the text.


land; Arnorsson, Sigurdsson, and Svavarsson, 1982; Arnorsson, Gunnlaugsson, andSvavarson, 1983) and assuming that the activity of Ca�2 was constrained by the absenceof grossular and paragonite in the assemblage (fig. 14). Resulting fluid compositions werebasic (pH 8.78-9.20) with activities of Ca�2 (log aCa

�2 � �5.59 to �4.99) that are 2 to 3.5orders of magnitude higher than the activity of Mg�2 (log aMg

�2 � �8.44 to �7.89).Calculated values of log [(aCa

�2)/(aMg�2)] agree well with values computed using compo-

sitions of Icelandic geothermal fluids at temperatures near 100°C (Arnorsson, Sigurds-son, and Svavarsson, 1982; Arnorsson, Gunnlaugsson, and Svavarson, 1983). Theabsence of calcite during Stage II requires that log fCO2 was less than �2.5 bars (fig. 14).

Porosity evolution.—Hydrogeochemical modeling of reactive fluid transport (Rose,1995; Manning, Ingebritsen, and Bird, 1993) and field observations (Manning and Bird,1991, 1995; Robert and Goffe, 1993; Bevins, Rowbotham, and Robinson, 1991; Schmidt,1990; Schmidt and Robinson, 1997) demonstrate the importance of porosity (�) andpermeability (k) in controlling mineral parageneses in low-grade metabasalts. In particu-lar, the magnitude of k influences the extent and distribution of chemical alterationduring hydrothermal and metamorphic processes (Norton, 1988; Rose, 1995). In regionsof large � and k, high fluid fluxes promote disequilibrium conditions that can influencethe locations of mineral isograds (Lasaga and Rye, 1993; Lasaga, 1989; Bolton, Lasaga,and Rye, 1999; Dewers and Ortoleva, 1990) and promote metastable mineral formation(Steefel and Van Cappellen, 1990; Steefel and Lasaga, 1994). Consequently, measure-ments of � and k are critical for modeling mineralogical alteration during fluid flow(Norton, 1988; Norton and Knapp, 1977). In lavas, these properties are controlled bycomplex sequences of pore forming events (for example, degassing of the lava to formvesicles during eruption, brittle deformation after extrusion, dissolution of primaryminerals during metasomatism) and reduction of � through mineral precipitations.Typical laboratory measurements of � (Norton and Knapp, 1977) in altered rocks reflectthe cumulative effects of these processes. Laboratory measurements on altered rockstypically underestimate k during active stages of hydrothermal and metamorphic flowsystems because much of the � is occluded by secondary minerals (Norton, 1988).Although laboratory measurements of initial � on unaltered rocks or outcrop measure-ments of pore abundance (such as those conducted in this study) can in part mitigatethese problems, it is clear that � and k change dramatically with time (generallydecreasing) during progressive infilling of pore spaces during mineralogical alteration(Ague, 1995). In addition, systematic decreases in k with depth in the crust (Manning andIngebritsen, 1999) can at least partially be attributed to mineral precipitation. Generationof secondary � associated with mineral dissolution during alteration (Ortoleva andothers, 1987) also complicates assessment of � and k at various stages in the alterationhistory of crustal rocks.

The data in tables 7 and 8 provide quantitative estimates of the amount of porositycreated during volcanism and brittle deformation (�*1 and �*2, respectively) and of theamount of �*1 destroyed through mineral precipitation. These data are summarized infigure 15, which depicts the evolution of � at Teigarhorn as a function of the three stagesin mineral paragenesis. Two main pore forming events influenced the lavas at Teigar-horn: creation of primary � during extrusion and solidification of the lavas and brittledeformation during dike emplacement. Vesiculation of the lavas resulted in a heteroge-neous distribution of �*1, with flow tops having as much as 22 percent �*1 and flow centersbeing essentially non-porous. Sporadic development of Stage I alteration decreased thevalue of � by �8 percent on average. Precipitation of Stage II C/S during burial resultedin a loss of an additional 40 percent of �. Stage II zeolites filled most of the remainingvolume of �*1. Development of Stage III alteration required addition of 3 to 11 percenttotal porosity (�*2) through brittle deformation. Quantitative measurements of the degreeof filling of �*2 by Stage III minerals were not conducted. However, outcrop observations


suggest that the percent of � occluded by mineral precipitation during Stage III wassmaller than that in Stage II as evidenced by the abundant open cavities remaining inbreccias and fractures that host museum-grade zeolite, quartz, and Iceland spar crystals.

The textures and extent of infilling of primary pore spaces during Stages I and IIsuggest that although � decreased during alteration, the matrix of the lavas wassufficiently permeable to permit transport of chemical components necessary for precipi-tation of observed mineral assemblages. As noted above, mass transfer must haveoccurred from the matrix of the lavas to primary pore spaces. This apparently was trueeven after pore walls were coated with Stage I silica and celadonite and Stage II C/S.With the exception of scoreaceous zones, where connections between pore spaces arereadily observable, much of the Stage I and II pore filling occurred in effectively isolatedvesicles. Vesicles hosting Stage II minerals were thus only partially sealed by precipita-tion of Stage I minerals and Stage II C/S. The nearly complete infilling of �*1 duringStage II was particularly acute because of the large molar volumes of zeolites andsmectites. Subsequent alteration was limited by the lack of pore space for storage andmigration of fluids, except where brittle deformation locally increased �, resulting inStage III alteration.

Although � was larger in flow tops at the onset of Stage II than in Stage III (table 7),fractures conducting fluid flow during Stage III are often laterally continuous in outcropfor as much as 10 m and associated with irregular porous metasomatized aureoles (figs.10, 11C and D). Thus, k during Stage III was probably much greater initially than inporous zones occupied by Stage I and II minerals, despite lower values of � (table 7).The extensive metasomatism producing the vuggy SiO2-rich zones near Stage III poreswas likely a consequence of high water/rock ratios possible in large, conductivefractures. Late precipitation of zeolites closed the fractures, resulting in a decrease of both� and k. The duration of convective fluid and heat transport during Stage III wasprobably limited by declining k due to fracture filling.

Fig. 15. Variation of bulk porosity (percent) in lavas at Teigarhorn as a function of paragenetic stage(based on the measurements in tables 7 and 8). Time elapsed after eruption increases to the right. Verticalbrackets illustrate the ranges in total primary and secondary porosity determined by outcrop and thin sectionmeasurements. The gray region illustrates changes in the range of bulk porosity values realized throughgeologic time, with the relative height and absolute position of the boundaries of this region illustrating thevariation in and magnitude of porosity.


CONCLUDING REMARKS

The mineral parageneses observed in eastern Iceland and other low-grade metaba-salt localities often result from multiple stages of alteration that can be difficult todistinguish. However, as shown in figure 13, the textural and mineralogical signatures oflow-grade metamorphism can be interpreted by relatively simple paragenetic schemesthat account for the tectonic history of the lavas. It is also readily apparent that mineralprecipitation and tectonic events accompanying the paragenetic stages depicted in figure13 profoundly influence the hydrogeologic properties of basaltic crust (fig. 15).

An example of the importance of petrotectonic interpretations of zeolite-faciesmineral parageneses is the ‘‘zone of abundant zeolites,’’ described by Walker (1960b) asthe deepest exposed mineral zone developed in non-olivine normative tholeiitic lavas ineastern Iceland. The assemblage described by Walker (1960b) in this zone (‘‘magnificentspecimens’’ of heulandite, stilbite, scolecite, epistilbite) is similar to the integratedmineralogy of Stage II and III alteration found at Teigarhorn, which Walker (1960b)cited as a type locality for the ‘‘zone of abundant zeolites.’’ Inspection of figures 3 and 4 inWalker (1960b) suggests that the vertical extent of the ‘‘zone of abundant zeolites’’ iscorrelated to dike distribution along the northern shore of Berufjordur just as thedistribution of Stage III alteration at Teigarhorn is associated with dikes. Based on thesecomparisons, it appears that the ‘‘zone of abundant zeolites’’ reported by Walker (1960b)represents hydrothermal aureoles around dike swarms and is not a signature of burialmetamorphism.

ACKNOWLEDGMENTS

The authors would like to thank Herbert Hjorleifsson for graciously granting usaccess to the outcrops at Teigarhorn, Elsa Thorey Eysteinsdottir, Gunnlaugur Fridbjarnar-son, and Obba, Oskar, and Berta Sandholt for friendship and hospitality, and the citizensof Djupivogur for accommodating us during our work at Teigarhorn. Anne Dyreborgassisted with the fluid inclusion analyses. This work was funded by the Shell and McGeeFunds at Stanford University, the Geological Society of America, the Icelandic StudentInnovation Fund, the Petroleum Research Fund (grant ACS-PRF 31742-AC2 to DKB),and graduate fellowships from the United States National Science Foundation andEnvironmental Protection Agency (to PSN). Discussions with L. Heister, A. Stefansson,and S. Jakobsson and reviews by J. Liou and P. Schiffman contributed significantly to ourinterpretations.

REFERENCES

Ague, J. J., 1995, Deep crustal growth of quartz, kyanite, and garnet into large-aperture, fluid-filled fractures,north-eastern Connecticut, USA: Journal of Metamorphic Geology, v. 13, p. 299-314.

Alberti, A., Pongiluppi, D., and Vezzalini, G., 1982, The Crystal Chemistry of Natrolite, Mesolite andScolecite: Neues Jarbuch fur Mineralogie, Abhandlungen, v. 143, p. 231-248.

Alt, J. C., 1999, Very low-grade hydrothermal metamorphism of basic igenous rocks, in Frey, M., andRobinson, D., editors, Low Grade Metamorphism: Oxford, Blackwell Science Ltd., p. 169-201.

Arnorsson, S., Gunnlaugsson, E., and Svavarsson, H., 1983, The chemistry of geothermal waters in Iceland; II,Mineral equilibria and independent variables controlling water compositions: Geochemica et Cosmo-chimica Acta, v. 47, p. 547-566.

Arnorsson, S., Sigurdsson, S., and Svavarsson, H., 1982, The chemistry of geothermal waters in Iceland; I,Calculation of aqueous speciation from 0 degrees to 370 degrees C: Geochemica et Cosmochimica Acta,v. 46, p. 1513-1532.

Bailey, S. W., 1988, Chlorites: structures and crystal chemistry, In Bailey, S. W., editor, Hydrous Phyllosili-cates: Mineralogical Society of America Reviews in Mineralogy, v. 19, p. 347-403.

Bauer, J., and Hrıchova, R., 1962, Ptilolite (mordenite) from Teigarhorn, Iceland: Sbornik Vysoke SkolyChemicko-Technologicke V Praze, Mineralogie, v. 6, p. 27-36.

Bellanca, A., 1940, Piezoelettricita e legge di geminazione della scolecite: Societa de Scienze Naturali EdEconomiche di Palermo. Bollettino, v. 22, p. 11-18.

Betz, V., 1981, Famous mineral localities; zeolites from Iceland and the Faeroes: The Mineralogical Record,v. 12, p. 5-26.

Bevins, R. E., Robinson, D., and Rowbotham, G., 1991, Compositional variations in mafic phyllosilicates fromregional low-grade metabasites and application of the chlorite geothermometer: Journal of MetamorphicGeology, v. 9, p. 711-721.


Bevins, R. E., Rowbotham, G., and Robinson, D., 1991, Zeolite to prehnite-pumpellyite facies metamorphismof the late Proterozoic Zig-Zag Dal Basalt Formation, eastern North Greenland: Lithios, v. 27, p. 155-165.

Blake, D. H., 1970, Geology of Alftafjordur volcano, a Tertiary volcanic centre in south-eastern Iceland:Visindafelag Islendinga [Rit], v. 2, p. 43-63.

Bolton, E. W., Lasaga, A. C., and Rye, D. M., 1999, Long-term flow/chemistry feedback in a porous mediumwith heterogeneous permeability: kinetic control of dissolution and precipitation: American Journal ofScience, v. 299, p. 1-68.

Buckley, H. A., Bevan, J. C., Brown, K. M., and Johnson, R., 1978, Glauconite and celadonite: Two separatemineral species: Mineralogical Magazine, v. 42, p. 373-382.

Cathelineau, M., 1988, Cation site occupancy in chlorites and illites as a function of temperature: ClayMinerals, v. 23, p. 471-485.

Cathelineau, M., and Nieva, D., 1985, A chlorite solid solution geothermometer. The Los Azulfres geothermalsystem (Mexico): Contributions to Mineralogy and Petrology, v. 91, p. 235-244.

Chayes, F., 1956, Petrographic model analysis: an elementary statistical appraisal: New York, John Wiley andSons, 113 p.

Christiansen, F. G., Boesen, A., Bojesen-Koefoed, J., Dallhoff, F., Dam, G., Neuhoff, P. S., Pedersen, A. K.,Pedersen, G. K., Stannius, L. S., and Zinck-Joergensen, K., 1999, Petroleum geological activities onshoreWest Greenland in 1997: Geology of Greenland Survey Bulletin, v. 180, p. 10-17.

Christensen, N. I., and Wilkens, R. H., 1982, Seismic properties, density, and composition of the Icelandiccrust near Reydarfjordur: Journal of Geophysical Research, v. 87, p. 6389-6395.

Dewers, T., and Ortoleva, P., 1990, Geochemical self-organization III: Mechano-chemical model of metamor-phic differentiation: American Journal of Science, v. 290, p. 473-521.

Flovenz, O. G., and Sæmundsson, K., 1993, Heat flow and geothermal processes in Iceland: Tectonophysics,v. 225, p. 123-138.

Flower, M. F. J., Pritchard, R. G., Brem, G., Cann, J. R., Delaney, J., Emmerman, R., Gibson, I. L., Oakley, P. J.,Robinson, P. T., and Schmincke, H.-U., 1982, Chemical stratigraphy, Iceland Research Drilling Project,Reydarfjordur, eastern Iceland: Journal of Geophysical Research, v. 87, p. 6489-6610.

Galli, R., and Rinaldi, R., 1974, The Crystal Chemistry of Epistilbites: American Mineralogist, v. 59,p. 1055-1061.

Ghiara, M. R., Franco, E., Petti, C., Stanzione, D., and Valentino, G.M., 1993, Hydrothermal interactionbetween basaltic glass, deionized water and seawater: Chemical Geology, v. 104, p. 125-138.

Gıslason, S. R., and Arnorsson, S., 1993, Dissolution of primary basaltic minerals in natural waters: saturationstate and kinetics: Chemical Geology, v. 105, p. 117-135.

Gıslason, S. R., Arnorsson, S., and Armannsson, H., 1996, Chemical weathering of basalt in Southwest Iceland;effects of runoff, age of rocks and vegetative/glacial cover: American Journal of Science, v. 296,p. 837-907.

Gıslason, S. R., and Eugster, H. P., 1987a, Meteoric water-basalt interactions. I: A laboratory study:Geochemica et Cosmochimica Acta, v. 51, p. 2827-2840.

————— 1987b, Meteroric water-basalt interactions. II: A field study in NE Iceland: Geochemica et Cosmo-chimica Acta, v. 51, p. 2841-2855.

Gıslason, S. R., Veblen, D. R., and Livi, K. J. T., 1993, Experimental meteoric water-basalt interactions:Characterization and interpretation of alteration products: Geochimica et Cosmochimica Acta, v. 57,p. 1459-1471.

Grant, J. A., 1982, The Isocon Diagram—A Simple Solution to Gresens’ Equation for Metasomatic Alteration:Economic Geology, v. 81, p. 1976-1982.

Gustafsson, L. E., 1992, Geology and petrography of the Dyrfjoll central volcano, eastern Iceland: BerlinerGeowissenschaftliche Abhandlungen, v. 138, p. 1-101.

Guven, N., 1988, Smectites, In Bailey, S. W., editor, Hydrous Phyllosilicates: Mineralogical Society of AmericaReviews in Mineralogy, v. 19, p. 497-559.

Heister, L. E., ms, 1997, The petrogenesis of Tertiary phonolitic volcanism Gronau West Nunatak, EastGreenland: M.S. thesis, Arizona State University, Tempe, Arizona, 12 p.

Holland, T., Baker, J., and Powell, R., 1998, Mixing properties and activity-composition relationships ofchlorites in the system MgO-FeO-Al2O3-SiO2-H2O: European Journal of Mineralogy, v. 10, p. 395-406.

Johannesson, H., and Sæmundsson, K., 1989, Geological Map of Iceland. 1:500000. Bedrock Geology:Reykjavik, Iceland, Icelandic Museum of Natural History and Iceland Geodetic Survey, 1 sheet.

Johnson, J. W., Oelkers, E. H., and Helgeson, H. C., 1992, SUPCRT92: Software package for calculating thestandard molar thermodynamic properties of minerals, gases, aqueous species, and reactions among themas functions of temperature and pressure: Computers in Geoscience, v. 18, p. 899-947.

Jørgensen, O., 1984, Zeolite zones in the basaltic lavas of the Faeroe Islands. In The deep drilling project1980-1981 in the Faeroe Islands: Annales Societatia Scientarum Faroensis, Supplementum, v. 9, p. 71-91.

Kristjansson, L., Gudmundsson, A., and Haraldsson, H., 1995, Stratigraphy and paleomagnetism of a3-km-thick Miocene lava pile in the Mjoifjordur area, eastern Iceland: Geologische Rundschau, v. 84,p. 813-830.

Kristmannsdottir, H., 1979, Alteration of basaltic rocks by hydrothermal activity at 100-300°C, in Mortland,M., and Farmer, V., editors, Developments in Sedimentology, v. 27: Amsterdam, Elsevier, p. 359-367.

Kristmannsdottir, H., and Tomasson, J., 1978, Zeolite zones in geothermal areas in Iceland, in Sand, L. B., andMumpton, F. A., editors, Natural Zeolites: Oxford, Pergamon Press Ltd., p. 277-284.

Lasaga, A. C., 1986, Metamorphic reaction rate laws and the development of isograds: MineralogicalMagazine, v. 50, p. 359-373.

————— 1989, Fluid flow and chemical reaction kinetics in metamorphic systems: a new simple model: Earthand Planetary Science Letters, v. 94, p. 417-424.


Lasaga, A. C., and Rye, D. M., 1993, Fluid flow and chemical reaction kinetics in metamorphic systems:American Journal of Science, v. 293, p. 361-404.

Liou, J. G., Seki, Y. S., Guillemette, R. N., and Sakai, H. U., 1985, Compositions and parageneses of secondaryminerals in the Onikobe geothermal system, Japan: Chemical Geology, v. 49, p. 1-20.

Manning, C. E., and Bird, D. K., 1991, Porosity evolution and fluid flow in the basalts of the Skaergaardmagma-hydrothermal system, East Greenland: American Journal of Science, v. 291, p. 201-257.

————— 1995, Porosity, permeability, and basalt metamorphism, in Schiffman, P., and Day, H. W., editors,Low-Grade Metamorphism of Mafic Rocks: Geological Society of America Special Paper 296, p. 123-140.

Manning, C. E., and Ingebritsen, S. E., 1999, Permeability of the continental crust: Implications of geothermaldata and metamorphic systems: Reviews in Geophysics, v. 37, p. 127-150.

Manning, C. E., Ingebritsen, S. E., and Bird, D. K., 1993, Missing mineral zones in contact metamorphosedbasalts: American Journal of Science, v. 293, p. 894-938.

Mehegan, J. M., Robinson, P. T., and Delaney, J. R., 1982, Secondary mineralization and hydrothermalalteration in the Reydarfjordur drill core, eastern Iceland: Journal of Geophysical Research, v. 87,p. 6511-6524.

Murata, K. J., Formoso, M. L. L., and Roisenberg, A., 1987, Distribution of zeolites in lavas of southeasternParana Basin, state of Rio Grande do Sul, Brazil: Journal of Geology, v. 95, p. 455-467.

Neuhoff, P. S., Watt, W. S., Bird, D. K., and Pedersen, A. K., 1997, Timing and Structural Relations of RegionalZeolite Zones in Basalts of the East Greenland Continental Margin: Geology, v. 25, p. 803-806.

Norton, D., 1988, Metasomatism and permeability: American Journal of Science, v. 288, p. 604-618.Norton, D., and Knapp, R., 1977, Transport phenomena in hydrothermal systems: The nature of porosity:

American Journal of Science, v. 277, p. 913-936.Odin, G. S., Desprairies, A., Fullagar, P. D., Bellon, H., Decarreau, A., Froehlich, F., and Zelvelder, M., 1988,

Nature and geological significance of celadonite, in Odin, G. S., editor, Green marine clays; ooliticironstone facies, verdine facies, glaucony facies and celadonite-bearing facies; a comparative study:Developments in Sedimentology, v. 45, p. 337-398.

Ortoleva, P., Chadam, J., Merino, E., and Sen, A., 1987, Geochemical self-organization II: The reactive-infiltration instability: American Journal of Science, v. 287, p. 1008-1040.

Oxburgh, E. R., and Agrell, S. O., 1982, Thermal conductivity and temperature structure of the Reydarfjordurborehole: Journal of Geophysical Research, v. 87, p. 6423-6428.

Palmason, G., 1973, Crustal rifting, and related thermo-mechanical processes in the lithosphere beneathIceland: Geologische Rundschau, v. 70, p. 244-260.

Passaglia, E., 1975, The crystal chemistry of mordenites: Contributions to Mineralogy and Petrology, v. 50,p. 65-77.

Passaglia, E., Galli, E., Leoni, L., and Rossi, G., 1978, The Crystal Chemistry of Stilbites and Stellerites: Bulletinde Mineralogie, v. 101, p. 368-375.

Ragnarsdottir, K. V., 1993, Dissolution kinetics of heulandite at pH 2-12 and 25°C: Geochimica et Cosmo-chimica Acta, v. 57, p. 2439-2449.

Reynolds, R. C. J., 1988, Interstratified clay minerals, in Brindley, G. W., and Brown, G., editors, CrystalStructures of Clay Minerals and Their X-ray Identification: London, Mineralogical Society, p. 249-304.

Robert, C., and Goffe, B., 1993, Zeolitization of basalts in subaqueous freshwater settings: Field observationsand experimental study: Geochimica et Cosmochimica Acta, v. 57, p. 3597-3612.

Robinson, D., and Bevins, R. E., 1999, Patterns of regional low-grade metamorphism in metabasites, in Frey,M., and Robinson, D., editors, Low Grade Metamorphism: Oxford, Blackwell Science Ltd., p. 143-168.

Robinson, D., Bevins, R. E., and Rowbotham, G., 1993, The characterization of mafic phyllosilicates inlow-grade metabasalts from eastern North Greenland: American Mineralogist, v. 78, p. 377-390.

Robinson, P. T., Mehegan, J., Gibson, I. L., and Schmincke, H.-U., 1982, Lithology and structure of thevolcanic sequence in eastern Iceland: Journal of Geophysical Research, v. 87, p. 6429-6436.

Rose, N. M., 1995, Geochemical consequences of fluid flow in porous basaltic crust containing permeabilitycontrasts: Geochemica et Cosmochimica Acta, v. 59, p. 4381-4392.

Sæmundsson, K., 1979, Outline of the geology of Iceland: Jokull, v. 29, p. 7-28.Schiffman, P., and Fridleifsson, G. O., 1991, The smectite-chlorite transition in drillhole NJ-15, Nesjavellir

geothermal field, Iceland: XRD, BSE and electron microprobe investigations: Journal of MetamorphicGeology, v. 9, p. 679-696.

Schmidt, S. Th., 1990, Alteration under conditions of burial metamorphism in the North Shore VolcanicGroup, Minnesota—Mineralogical and geochemical zonation: Heidelberger Geowissenschaftliche Abhand-lungen, v. 41, 309 pp.

Schmidt, S. Th., and Robinson, D., 1997, Metamorphic grade and porosity and permeability controls on maficphyllosilicate distributions in a regional zeolite to greenschist facies transition of the North Shore VolcanicGroup, Minnesota: Geological Society of America Bulletin, v. 109, p. 683-697.

Seki, Y., Liou, J. G., Guillemette, R., Sakai, H., Oki, Y., Hirano, T., and Onuki, H., 1983, Investigation ofgeothermal systems in Japan. I. Onikobe Geothermal Area: Saitama University 3, Hydroscience andGeotechnology Laboratory, Memoir, 206 pp.

Shau, Y.-H., and Peacor, D. R., 1992, Phyllosilicates in hydrothermally altered basalts from DSDP Hole 504B,leg 83_A TEM and AEM study: Contributions to Mineralogy and Petrology, v. 112, p. 119-133.

Slaughter, M., and Kane, W. T., 1969, The crystal structure of a disordered epistilbite: Zeitschrift furKristallographie, v. 130, p. 68-87.

Steefel, C. I., and Lasaga, A. C., 1994, A coupled model for transport of multiple chemical species and kineticprecipitation/dissolution reactions with applications to reactive flow in single phase hydrothermalsystems: American Journal of Science, v. 294, p. 529-592.


Steefel, C. I., and Van Cappellen, P., 1990, A new kinetic approach to modeling water-rock interaction: therole of nucleation, precursors, and Ostwald ripening: Geochimica et Cosmochimica Acta, v. 54, p. 2657-2677.

Sukheswala, R. N., Avasia, R. K., and Gangopadhyay, M., 1974, Zeolites and associated secondary minerals inthe Deccan Traps of Western India: Mineralogical Magazine, v. 39, p. 658-671.

Tingle, T. N., Neuhoff, P. S., Ostergren, J. D., Jones, R. E., and Donovan, J. J., 1996, The effect of ‘missing’(unanalyzed) oxygen on quantitative electron microprobe microanalysis of hydrous silicate and oxideminerals: Geological Society of America, Abstracts with Programs, v. 28, p. A-212.

Walker, G. P. L., 1951, The amygdale minerals in the Tertiary lavas of Ireland. I. The distribution of chabazitehabits and zeolites in the Garron plateau area, County Antrim: Mineralogy Magazine, v. 29, p. 773-791.

————— 1959, Geology of the Reyoarfjorour area, eastern Iceland. The Quarterly Journal of the GeologicalSociety of London, v. 114, p. 367-393.

————— 1960a, The amygdale minerals of the Tertiary lavas of Ireland. III. Regional distribution: MineralogicalMagazine, v. 32, p. 515-528.

————— 1960b, Zeolite zones and dike distribution in relation to the structure of the basalts of eastern Iceland:Journal of Geology, v. 68, p. 515-528.

————— 1962, Tertiary welded tuffs in eastern Iceland: The Quarterly Journal of the Geological Society ofLondon, v. 118, p. 275-293.

————— 1963, The Breiddalur central volcano, eastern Iceland [with discussion]: The Quarterly Journal of theGeological Society of London, v. 119, p. 29-63.

————— 1974, The structure of eastern Iceland: in Kristjansson, l., editor, Geodynamics of Iceland and the NorthAtlantic Area: Reykjavik, Iceland, NATO Advanced Study Institute, p. 177-188.

Watkins, N. D., and Walker, G. P., L., 1977, Magnetostratigraphy of eastern Iceland: American Journal ofScience, v. 277, p. 513-584.

Wolery, T. J., 1991, EQ3NR. A computer program for geochemical aqueous speciation solubility calculations:Theoretical manual, user’s guide and related documentation (Version 7.0). Livermore, California,Lawrence Livermore National Laboratory report UCRL-MA-110662 Pt III.

Wood, D. A., 1976, Elemental Mobility during Zeolite Facies Metamorphism of the Tertiary Basalts of EasternIceland: Contributions to Mineralogy and Petrology, v. 55, p. 241-254.

————— 1977, Major and trace element variations in the Tertiary lavas of eastern Iceland and their significancewith respect to the Iceland geochemical anomaly: Journal of Petrology, v. 19, p. 393-436.


Documents

porosity evolution and mineral paragenesis during low-grade